or chlorite in the fossil hydrothermal system of Saint Martin (Lesser Antilles)

Journal o f Volcanology and Geothermal Research, 51 (1992) 95-114 Elsevier Science Publishers B.V., Amsterdam

95

Chemical variations in assemblages including epidote and/or chlorite in the fossil hydrothermal system of Saint Martin (Lesser Antilles) D. BeauforP, P. Patrier~, A. M e u n i e l~ a n d M . M . O t t a v i a n i b ~Laboratoire de pbtrologie des altbration hydrothermales, Universitb de Poitiers, U.A. 721 C.N.R.S., 40 avenue du recteur Pineau, 86022 Poitiers cedex, France bSciences de la terre, Universitb de Corse, Avenue J. Nicoli, B.P. 24, 20250 Corte, France (Received April 3, 1989; revised and accepted November 20, 1991 )

ABSTRACT Beaufort, D., Patrier, P., Meunier, A. and Ottaviani, M.M., 1992. Chemical variations in assemblages including epidote and/or chlorite in the fossil hydrothermal system of Saint Martin (Lesser Antilles). J. Volcanol. Geotherm. Res., 51: 95114. Epidote and/or chlorite are common minerals in the roots of the fossil geothermal system of Saint Martin (Lesser Antilles). They appear in four distinct assemblages: ( 1) epidote + actinolite+ quartz + magnetite near the contact between the tuffaceous host rocks (andesitic modal composition) and the quartz-diorite intrusion of Philipsburg; (2) epidote+chlorite+quartz in host rocks as far as a lateral distance of about 3 km from the intrusion; (3) epidote + chlorite + haematite + quartz locally in iron and manganese rich host rocks; (4) chlorite + phenglte + magnetite appearing as late sealing of porosity in fracture-controlled quartz veins with strongly phengitized wall rocks. All these assemblages constitute a large alteration grading from propylitic alteration to thermal metamorphism (actinolite-bearing assemblage). Detailed microprobe studies of epidotes replacing plagioclases and of chlorites replacing glass and mafic minerals reveal notable compositional variations which have been studied with respect to temperature paleogradients (estimated from fluid inclusions study), distance from the thermal source and fo2 conditions. The mean Ps + Pm [ 100 X (Fe a+ + Mn 3+ )/ (AP + + Fe a+ + Mn a+ ) ] ofepidotes vary from 21 in the presence of magnetite near the intrusion to 32 in haematite-bearing iron and/or manganese volcanic and sedimentary formations. The intra-grain chemical scattering of epidotes increases with increasing distance of the pluton and decreasing temperature of crystallization. All the chlorites coexisting with epidote are Mg-rich ()tEe< 0.50). Their main compositional variation consists in a significant enrichment in magnesium (toward the chlinochlore end member) in presence of haematite. The intra-grain chemical scattering of chlorite (expressed by the aluminium content in the structural formula) increases with increasing distance of the pluton and decreasing temperature of crystallization. Chlorites associated with phengite and magnetite in vein alteration are Fe- and Al-rich. The M6ssbauer spectra indicate that the Fe3+ content of chlorite varies between 25 and 32% of total Fe in the presence of epidote; the higher content being attained in the presence of haematite. The Fe3+ content of chlorite associated with magnetite and phengite is 16% of total Fe. The compositional variations of epidote and/or chlorite of the four distinct assemblages observed at Saint Martin result from the combined effects offo2, temperature, and time of heating. The effect offo2 is particularly perceptible in the control of the epidote Ps content, of the chlorite XFe ratio of Fe 3+ distribution between coexisting epidotes and chlorites. Despite the fact that it may be partially canceled out by the effect offo2, the variation of compositional ranges of both epidotes and chlorites, which increases toward the outer part of the geothermal system in response to the combination of decreasing temperatures and decreasing time of heating of the rocks, suggests that chemical equilibrium has not been attained in the assemblages bearing epidotes and chlorites.

Correspondence to: D. Beaufort, Laboratoire de p6trologie des alt6ration hydrothermales, Universit6 de Poitiers, U.A. 721 C.N.R.S., 40 avenue du recteur Pineau, 86022 Poitiers Cedex, France.

0377-0273/92/$05.00 © 1992 Elsevier Science Publishers B.V. All fights reserved.

96

Introduction Epidote and chlorite are major rock-forming minerals in diverse geological environments subjected to moderate temperature conditions (200-350 ° C). Being indicators of greenschist facies in regional metamorphism, epidote and chlorite are also very common in contact metamorphism and hydrothermal alteration. In geothermal systems, coexisting epidote and chlorite appear near the bottom of the deep drilling (Tomasson and Kxistmannsdottir, 1972; Elders et al., 1979; Cavaretta et al., 1982; Cathelineau et al., 1985; Huelen and Nielson, 1986). This assemblage occurs in great volumes of altered rocks, well known as propylitic alteration (Creasey, 1959; Lowell and Guilbert, 1970; Beane, 1982). In various chemical systems, phase relations including solid solutions have been investigated independentlyfor epidote (Strens, 1965; Holdaway, 1972; Bird and Helgeson, 1980, 1981; Liou et al., 1983) and for chlorite (Nelson and Roy, 1958; Fawcett and Yoder, 1966; Velde, 1973; Mac Onie et al., 1975; Fleming and Fawcett, 1976; and Bryndzia and Scott, 1987, among an abundant literature). During the last few years, several works focused on the epidote-chlorite hydrothermal assemblage in order to provide tools for economic geology and for theoretical aspects of alteration petrology. The compositional trends of these minerals have been investigated as possible indicators of paleoconditions (Exley, 1982, Cavaretta et al., 1982 ). In this respect, a few authors proposed some applications of the compositional variations of hydrothermal chlorites to geotherrnometry (Mac Dowell and Elders, 1980; Cathelineau and Nieva, 1985; Walshe, 1986). Nevertheless, these geothermometers would be more reliable if they did not ignore the compositional variations of the other silicates coprecipitated with the chlorites. Only a few studies focused on the geochemical variations of coexisting epidote and chlorite in geothermal fields. At the present day, a

D. BEAUFORT ET AL.

general interpretation of the chemical trends observed in coexisting epidote and chlorite from geothermal systems remains tentative because of complex solid solution models for each mineral [Stoessel (1984) and Walshe (1986) proposed a six-component model to interpret the chemical variations in chlorite], and because of variable and complex zonal relationships exhibited by the epidotes (Bird et al., 1984; Shikazono, 1984). This communication focuses on the concomitant compositional variations of epidotes and chlorites in a fossil geothermal system generated by the intrusion of a quartz-diorite pluton into the volcanic and sedimentary formations of the island of Saint Martin (Lesser Antilles ). The objectives of the present study were: ( 1) to clarify the phase relations in the different assemblages including epidote and/or chlorite; (2) to determine the compositional variations of epidote and chlorite in the system; and (3) to outline and discuss the combined influence of temperature and oxygen fugacity on these chemical variations.

General petrography and hydrothermal alteration The geology and the lateral distribution of the extensive halo of hydrothermal alterations surrounding the quartz-diorite pluton of Philipsburg (emplaced during the Early Oligocene) have been described by Beaufort et al. (1990). The gross stratigraphy of the Eocene volcanic and sedimentary host formations consists of an upper fine-grained hyaloclastite sequence, a middle sequence of andesite lava flows and hyaloclastite and a lower sequence of alternating marls, limestones and minor hyaloclastites. In spite of weak structural deformation, these sequences were altered as far as 4 km from the quartz-diorite intrusion; they show a typical alteration zoning pattern, locally crosscut by quartz veins with strongly sericitized wall rocks identified as "phyllic veins". The petrological characteristics of the altered

CHEMICALVARIATIONSIN ASSEMBLAGESINCLUDINGEP1DOTEAND/ORCHLORITE

rocks and the fluid inclusion data argue that the alteration system, which is accessible at the present day, corresponds to the roots of a large geothermal circulation system. Figure 1 summarizes the lateral distribution of the hydrothermal assemblages in the Eocene host formations of Saint Martin. Secondary epidote and chlorite crystallized in the inner part of the system and may be continuously observed as far as about 3 km from the pluton. Five assemblages have been distinguished: ( 1 ) An epidote + actinolite + quartz + magnetite assemblage appears at the periphery of the pluton and in the surrounding host rocks Ikm

97

as far as 300 m from their contact; it corresponds to the thermal aureole. (2) Outside the above-mentioned assemblage an epidote÷chlorite+quartz assemblage replaces the volcanic and sedimentary rocks as far as 3 km from the pluton. The alteration of Fe-rich tuffaceous rocks including several Fe-Mn rich cherty layers (between 1.2 and 1.7 km from the intrusion) gives way to: (3) An epidote + chlorite + haematite + quartz assemblage; these rocks in which the actinolite-bearing assemblage crystallized were structurally modified showing a typical polygonal structure (hornfels facies). The primary

," Kool Bay

Fort HII"I

-"

4

Quartz diorite

I

lEoceneformations

r I ' ;Miocenelimestones

TEMPERATURE l°C) . |Sl

/ I

lse

O

"N

i

i

:

DISTANCEFROMTHE INTRUSION(kin) Fig. 1. Simplified lateral cross section (Fort Hill-Kool Bay) of the fossil geothermal system of Saint Martin. The paleotemperature profile was estimated from the study of fluid inclusions (Beaufort et al., 1990 l=epidote +actinolite_+ magnetite assemblage; 2 = epidote + chlorite assemblage; 3 = epidote + chlorite + haematite assemblage; 4 = chlorite _+phengite_+ magnetite in phyllic veins; 5 = interlayered chlorite/smectite_+ illite/smectite mixed-layers + calcite assemblage (clay-carbonate outer zone).

98

13. BEAUFORT ET AL.

structures of the rocks in which the chloriteepidote bearing assemblages crystallized were preserved (propylitic facies). These propylitic rocks show pervasive replacement of plagioclase by epidote + quartz + residual albite and replacement of mafic minerals and glass by chlorite + quartz + haematite +__minor epidote. (4) The chlorite + magnetite + phengite assemblage is restricted to the wall rock of"phyllic veins". Small grains of magnetite appear in chlorite patches. Petrographic considerations suggest that these chlorites sealed the residual porosity during the closing stage of the phyllic vein alteration. (5) At Saint Martin, the outer lateral part of the hydrothermal system consists of a calcite +chlorite/smectite_+ illite/smectite mixedlayer minerals +_haematite assemblage.

Sampling and method of investigation

The material of this study consists of 21 samples distributed over a lateral distance of 3 km (Fig. 2 ). Among these samples, three correspond to the epidote + actinolite_ magnetite+quartz assemblage, one to the epidote + actinolite + chlorite assemblage, seven to the epidote+chlorite+quartz assemblage, six to the epidote + haematite + chlorite + quartz assemblage and four samples correspond to the chlorite +_phengite + magnetite related to veins. All these samples were prepared as polished thin sections and mineral analyses were made on a CAMECAMS 46 probe equipped with an energy dispersive system E.G.G. ORTEC. Analyses were performed with a low-energy electron beam 0.001 pA-15 kV) with a spot size of 5 #m; the counting time was 100 seconds. Elements routinely analysed were Na, Mg, A1, Si,

MARIGOT

/

/

\ "sHIL

/

ERG

i 56

\

KOOL BAY

I 14 10

25

I 26

20 27

29 30 t9

J

Fig. 2. Location of the samples analysed in the Fort Hill-Kool Bay area.

KLE IN BAY

FORT ~AMSTI

CHEMICALVARIATIONSIN ASSEMBLAGESINCLUDING EPIDOTEAND/OR CHLORITE TABLE 1 Mean microprobe analyses and structural formulas of epidotes replacing plagioclase and biotite in sample SM33 Sample:

SM33

Allogenicparents: An. in aver.:

Plagioclases 24

Biotites 8

SiO2 A1203 MgO TiO 2 Mn203 CaO Na20 K20 Total

37.66 24.19 12.02 0.15 0.08 0.09 23.80 97.99

37.70 23.43 13.65 0.19 0.09 0.41 22.61 98.08

Si AI Fe a+ Mg Ti Mn 3+ Ca Na K Ps+Pm

5.97 + 0.03 4.53 + 0.06 1.43_+ 0.08 0.03 0.01 0.01 4.05 + 0.07 25+ 1

5.99_+ 0.03 4.40 + 0.02 1.63 + 0.05 0.05 0.01 0.05 3.86 _+0.13 28+ 1

Fe203

Structural analyses have been calculated on the basis of 25 oxygens; total iron has been considered as Fe 3+; Ps + P m = 100 (Fe3+ + Mn3+)/(Fe3+ + Mn3+ + A P ÷ ) ; +=standard deviation.

K, Ca Ti, Mn and Fe. The system was calibrated with a variety of synthetic oxide and natural silicate standards; matrix corrections were made with a Z.A.F. computing program. The reproductibility of standard analyses is 2% for Na and 1% for the other elements. In each sample, the presence of chlorite was verified by X-ray diffraction (Philips PW 1730 diffractometer, Co anticathode, 40 kV-40 mA), in order to avoid the presence of chlorite/smectite interstratified minerals with low range of smectite layers (so called "swelling chlorite"). These minerals have been identified in the peripheral part of the system (Beaufort et al., 1990). Optically similar to chlorite, they were formed at a slightly lower temperature. X-ray patterns differ from "true" chlorite by a low range of expandability after ethylene

99

glycol saturation and by a stronger (001) basal reflection. Chemically they differ from "true" chlorite by having higher silica, calcium and/ or alkali content and lower octahedral occupancy. They are commonly mistaken for chlorite and led to erroneous interpretations of chemical variations observed in "chlorites" crystallized for temperature conditions lower than 200 ° C (Bettison and Shiffman, 1988 ). Due to both inter-grain and intra-grain chemical variations, the determination of compositional trends of epidote required analysis at a large numbers of points in each sample. Inter-grain compositional variations are often the result of local equilibrium systems in which the chemistry of the newly formed minerals was controlled by the chemistry of their allogenic parents (Beaufort and Meunier, 1983). Thus, epidotes replacing marie minerals contain higher Fe than those replacing feldspars (Table I ). A minimum of twenty points have been analyzed in epidote per sample. The authors have minimized the effect of local chemical equilibria by considering only epidotes which replaced plagioclases (they constitute the major host for crystallization of epidote). The microprobe analyses of epidotes, chlorites and coexisting silicates are presented as mean values. The standard deviations for major elements illustrate The chemical heterogeneity. The structural formulas have been calculated on the basis of 25 oxygens for epidotes, 23 oxygens for actinolites, 28 oxygens for chlorites and 22 oxygens for phengites. Bulk-rock chemical analyses were obtained with an atomic absorption spectrometer (Perkin Elmer 107 ) after fusion of samples and dissolution by HNO3 digestion. Mtissbauer spectra were obtained from finely ground powders of four chlorite concentrates at room temperature. They were obtained on an Elsint A.M.E. 30 spectrometer (Mrtallurgie Physique laboratory, Poitiers University). A 57Co in Rh source of nominal 25 mCi was used. Signals were recorded on a multichannel ana-

100

D. B E A U F O R T ET AL.

35

Ii I iI+I I

30

+ O+-

25

20 !

~

I

i

o

DISTANCE FROM THE INTRUSION (kin) Fig. 3. G r a p h o f m e a n P s + P m in e p i d o t e v e r s u s lateral d i s t a n c e f r o m t h e q u a r t z - d i o r i t e p l u t o n . Vertical b a r s c o r r e s p o n d to s t a n d a r d d e v i a t i o n . Solid circle = c h l o r i t e a s s o c i a t e d w i t h e p i d o t e w i t h o u t i r o n o x i d e m i n e r a l ; solid t r i a n g l e = c h l o r i t e associated with epidote + haematite; solid square = chlorite associated with phengite + magnetite.

TABLE 2

TABLE 3

Mean microprobe analyses and structural formulas of epidotes associated with actinolite _+magnetite near the pluton

Mean microprobe analyses and structural formulas of actinolites associated with epidote -+magnetite near the pluton

Samples: An. in aver.:

SM 107 30

SM43A 31

SM40B 28

Samples: SM33 An. in 27

SiO2 AI203 Fe203 MgO TiO2 Mn203 CaO Na20 K20 Total

37.71 25.38 10.83 0.25 0.03 0.27 22.41 96.88

37.63 25.40 9.74 0.15 0.10 0.49 23.83 0.01 97.35

37.75 25.70 9.51 0.10 0.25 1.29 23.41 -

Si AI Fe 3+ Mg Ti Mn 3+ Ca Na K Ps+Pm

6.00_+ 0.03 4.80 _+0.06 1.25_+0.07 0.07 0.03 3.81 -+0.06 22+2

5.97 _+0.05 4.75 _+0.07 1.16_+0.08 0.03 0.01 0.06 4.05 _+0.12 22_+ 1

SM107 10

SM43A 4

SM40B 7

57.97 1.94 12.42 10.50 0.07 1.60 11.80 0.81 0.03 97.14

55.61 1.00 15.83 14.20 0.21 0,25 10.41 0,20 0.20 97.91

55.26 1.12 14.59 14.36 0.20 0.16 11.36 0.35 0.09 97.49

aver.:

98.01 5.95 -+ 0.05 4,77 _+0.07 1.13+_0.07 0.02 0.03 0.14 3.95 _+0.13

21-+ 1

Structural analyses have been calculated on the basis o f 25 oxygens; total iron has been considered as Fe3+; P s + P m = 1 0 0 (Fe 3+ + M n 3+ ) / ( F e 3+ + Mn 3÷ +AI 3+ ); _+ = s t a n d a r d deviation.

SiO2 A1203 FeO MgO TiO2 MnO CaO Na20 K20 Total Si Al Fe 2+ Mg Ti Mn 2÷ Ca K

56.05 1.24 3,39 20.42 0.08 5.26 11,58 0.08 0.10 98.20 7.82_+0.08 0.20-+0.06 0.40-+0.04 4.25_+0.06 0.02 0.62_+0.08 1.73_+0.03 0.01

7.94_+0.04 0.37_+0.04 1.67_+0.08 2.51-+0.11 0.01 0.22_+0.05 2.03_+0.04 0,01

8.02_+0.05 0.17_+0.03 1.91_+0.06 3.05-+0.06 0.03 0.03 1.61 +0.03 0,04

7.99__+0.06 0.19_+0.04 1.76+0.12 3.09_+0.10 0.02 0.02 1.76_+0.06 0.02

Structural analyses have been calculated on the basis of 23 oxygens; total iron has been considered as Fe2+; + = s t a n d a r d deviation.

CHEMICALVARIATIONSIN ASSEMBLAGESINCLUDINGEPIDOTE AND/OR CHLORITE

lyser. Isomer shifts were calculated vs. Fe metal. Deconvolutions, based on the leastsquares fitting procedure, assumed lorentzian line shapes. Each absorption doublet was characterized by the isomer shift, quadrupole splitting, peak intensity and line width.

101

piemontite

(Pm) content = 100, Ca4Mn6 Consequently, the chemical trend of the epidotes may be expressed as the variation of Ps + Pm content (Fig. 3 ): ( 1 ) The coarse-grained epidotes (length > 1 m m ) coexisting with iron-bearing actinolite and minor magnetite near the pluton (Table 2) are aluminium-rich. Their mean P s + P m ratio varies from 21 to 22 with standard deviation of 1 to 2. The most aluminium-rich point analysed is Ps + Pm = 18. These variations are due to intra-grain chemical zoning optically characterized by several abrupt variations of natural colour (pale yellow to pale green ). The analyses of associated iron-bearing actinolite (Table 3) indicate mean FeO contents (arbitrarily considered as Fe 2+ ) ranging from 12.42 to 15.83 wt.% and m1203contents ranging from 1 to 1.94 wt.%. The small excess of silicon in the structural formula is due to the presence of 5 i 6 0 2 4 ( O H ) 2 ]"

Microprobe analyses of epidotes All the epidotes (438 analyses) from the fossil system of Saint Martin are iron-rich. Most of them contain minor quantities of manganese; however, in Fe-Mn rich cherts, the Mn content in epidote may be substantial (values up to 10 wt.% M n 2 0 3 have been measured in sample SM19). The main chemical variations of epidotes are due to A13+ = F e 3+ substitution [expressed as pistacite (Ps) content = 100. C a 4 F e 6 S i 6 0 2 4 ( O H ) 2 ] and a minor part to A13+ = Mn 3+ substitution [expressed as TABLE 4

Mean microprobe analyses and structural formulas of epidotes associated with chlorite + quartz in absence of iron oxide mineral in volcanic and sedimentary host rocks Samples: An. in aver.:

SM56 38

SM04 30

SM09 29

SMI0 16

SM14 27

SMI7 29

SM21 17

SM33 24

SiO 2 A1203 Fe203 MgO TiO2 Mn203 CaO Na20 K20 Total

36.71 24.88 10.25 0.02 0.09 0.17 23.40 0.05 95.57

37.52 23.18 14.55 0.07 . 0.14 23.08 . 0.01 98.55

37.10 22.39 14.78 .

37.17 22.93 13.81 . . 0.17 22.73

37.19 23.77 12.64

37.69 23.61 12.02 0.02

0.35 23.08

0.30 23.24

37.66 24.19 12.02 0.15 0.08 0.09 23.80

96.81

38.06 23.92 13.09 . . 0.13 23.31 . . 98.51

97.84

97.99

Si A! Fe 3+ Mg Ti Mn 3+ Ca Na K Ps+Pm

5.94_+0.04 5.96_+0.10 5.99_+0.06 4.75-+0.19 4.34_+0.23 4.26-+0.24 1.25_+0.24 1.74_+0.27 1.70_+0.24 0.01 0.02 . . 0.01 . . . 0.02 0.02 0.05 4.06 3.93+0.04 3.86-+0.1 0.02 . . . . . . . 21+4 29_+4 30_+4

5.99_+0.03 4.36-+0.18 1.68_+0.19 . . 0.02 3.93+0.08 . . 28_+2

6.03_+0.07 4.34_+0.20 1.62_+0.23 . . . . 0.01 3.99-+0.07 . . . . 27-+3

.

. .

0.38 22.34 .

. . 96.99

. .

.

.

.

. . 97.03

.

5.97_+0.05 4.50_+0.17 1.53_+0.16

5.98_+0.04 4.54_+0.10 1.43+0.08

0.04 3.97_+0.07 . . 26-+2

0.04 3.95_+0.02

25_+2

5.97-+0.03 4.53_+0.06 1.43-+0.08 0.03 0.01 0.01 4.05_+0.07

25+1

Structural analyses have been calculated on the basis of 25 oxygens; total iron has been considered as Fe 3+ and total Mn as Mn 3+; P s + P m = I00 (Fe 3+ + M n 3+ ) / ( F e 3+ + M n 3+ +AI 3+ ); + :-standard deviation.

102

D. BEAUFORT ET AL.

TABLE 5 Mean microprobe analyses and structural formulas o f epidotes associated with chlorite + haematite in volcanic and sedimentary host rocks Samples: An. in aver.:

SM23 15

SM26 16

SM27 26

SM29 24

SM30 27

SM 19 21

SiO2 A1203 Fe203 MgO TiO2 Mn:O3 CaO Na:O KzO Total

37.21 22.34 15.11 0.38 22.94 . 97.98

37.53 23.30 14.70 0.10 0.09 0.48 22.26

37.45 21.76 16.21 0.17 0.01 0.22 22.12 . 0.04 97.98

37.19 22.62 14.90 0.23 23.01

36.38 21.85 15.56 0.03 0.27 22.48

37.42 21.86 1 1.92 0.21 0.10 2.76 23.30

0.02 96.59

0.01 97.53

Si AI Fe 3÷ Mg Ti Mn 3÷ Ca Na K Ps+Pm

5.96 + 0.03 4.22+0.18 1.82 + 0.21 0.05 3.94_ 0.06 . 31 _+2

.

. 0.03 98.49 5.96 + 0.04 4.36+0.11 1.76 + 0.13 0.02 0.01 0.06 3.78 + 0.06 .

. 0.01 30+2

.

. 97.95

6.00 + 0.05 4.11 + 0 . 2 0 1.96_+ 0.25 0.04 0.03 3.80 + 0.07 . 0.01 33_+3

5.95 + 0.02 4.27+0.07 1.80_+ 0.04

0.03 3.96 + 0.04 . 30_+ 1

5.93 + 0.03 4.20+0.05 1.91 _+0.05 0.01 0.03 3.92_+ 0.08 32+ 1

6.03 +_0.06 4.15 +0.19 1.44 + 0.21 0.05 0.01 0.30 4.02_+ 0.17 30+3

Structural analyses have been calculated on the basis of 25 oxygens; total iron has been considered as Fe 3+ and total Mn as Mn 3+; P s + P m = 100 (Fe 3+ + Mn 3+ ) / ( F e 3+ + Mn 3÷ + A P + ); + = s t a n d a r d deviation.

a part of the iron in ferric state in these minerals (Deer et al., 1962). (2) The compositions of epidotes associated with chlorite and quartz (Table 4) indicate a variation of the mean Ps + P m from 21 to 30 over a distance ranging from 500 m to 3 km outward from the pluton. Even if the mean P s + P m value of the sample SM56 is low (21), the epidotes with the highest values of mean Ps + P m seem rather located toward the external part of the system. These chemical variations are accompanied by a gradual increase of the standard deviation of Ps + P m from 1 to 4 and by a significant decrease of grain size (length decreasing from 300 to 50 # m ) toward the external part of the geothermal system. The increasing of the epidotes chemical variability as a function of the distance to the thermal source is mainly due to intra-grain chemical

heterogeneity. (3) The composition of greenish epidotes associated with chlorite + quartz + haematite (Table 5) indicates mean P s + P m ranging from 30 to 33. Mn-rich epidotes (mean value 2.76 Mn203 wt.% in sample SM 19 ) crystallized in Mn-rich iron cherts interlayered within the fine-grained hyaloclastite sequence. Epidote with a mean composition of Ps + Pm = 32 to 33 crystallized in the most haematite-rich samples (SM27 and SM30). The most iron-rich point analysed in epidote (measured in sample SM30) is P s + P m = 36. In the presence of haematite, the standard deviation of P s + P m ranges from 1 to 3. No clear correlation may be established between the chemical heterogeneity of grains and their distance from the pluton as was observed for the epidote + quartz + chlorite assemblage.

103

CHEMICAL VARIATIONS IN ASSEMBLAGESINCLUDING EPIDOTE AND/OR CHLORITE

1.0C

.80

[] .60.

O rid

BRUNSVIGITE

+

#.

+

•

.40.

ee

.20_ CLINOCHLORE

0

I

5 20

5.40

I

5.60

I

5 80

I

6.00

6 20

Silicon Fig. 4. Plot of chlorite mean analyses in Fe/(Fe+Mg+Mn) versus Si diagram (Foster, 1962). Solid circle=chlorite associated with epidote without iron oxide mineral; solid triangle= chlorite associated with epidote+ haematite; open square= chloriteassociatedwith phengite+ magnetite.Total Fe is consideredas Fe2+.

Microprobe analyses of chlorites All the chlorites identified by X-ray diffraction in the geothermal system of Saint Martin have the IIb polytype. The compositional data listed in Tables 6-8 summarize the results of 282 points analysed within eighteen samples. The structural formulas were calculated on the basis of 28 oxygens, assuming the total iron as Fe 2+. According to their paragenetic association, the most significant compositional shifts of chlorites concern XF, ratio ( F e + M n / F e + M n + M g ) coupled with Si and A1 distribution in the structure. The graphic representation adopted from Foster (1962) seems appropriate to illustrate these major chemical trends (Fig. 4) and the variations of Xr~ of both chlorite and bulk rock of weakly fractured rocks with the distance from the intrusion are presented in Figure 5a. ( 1 ) The composition of chlorites associated

with quartz and epidote ( P s + P m = 2 1 to 30) corresponds to Mg brunsvigite according to the classification of Foster ( 1962). Their mean XFe value ( F e + M n / F e + M g + M n ) ranges from 0.31 to 0.50 (Table 6). The intersample XFe variations seem partly correlated to the bulk rock chemistry (Fig. 5b). The standard deviations on XF~ vary between 0.01 to 0.03. These variations mainly result from intergrain analyses. (2) In presence of haematite, epidote (Ps + Pm = 30) and quartz, the composition of chlorite ranges over the Mg-rich brunsvigite and clinochlore compositional field. The mean XF~ varies between 0.34 and 0.08 (Table 7) and the standard deviation of XF~ varies between 0.01 and 0.02. Compared to the previous described brunsvigites, the chlorites associated with haematite are richer in silicon. There is a negative correlation between the total iron content of the bulk rock and the mean XF~ in associated chlorites (Fig. 5c). The most Mg-

104

D. BEAUFORTET AL.

(a)

1.0,

.8 .7

.6 .5 ,4 .3 .2

£

•

.........

.-.o,

,,' " O ~ a

"''---

O-_.#J

/

-

BUI,K-ROCK

"'lit

CItLORITE

.1 0

DISTANCE FROM THE INTRUSION (km)

.70-

1.00_

(c)

(b)

.80

.6o+

+

+

+ +

+

.50_

.40

I

.30

.40

I

.50

Fe + Mn ................ Fe + Mg +Mn

I

.60

chlorite

.60

A0

I

0

I

.20

I

I

.40

Fc + Mn ................ Fe + Mg +Mn

I

I

.60

chlorite

Fig. 5. G r a p h o f m e a n Xre in chlorite a n d in bulk rock from weakly fractured rocks versus lateral distance from the quartzdiorite pluton ( a ) . Plot o f chlorite XFe versus bulk-rock XFe for chlorite + epidote ( b ) or chlorite + epidote + h a e m a t i t e bearing samples (c). Total Fe is considered as Fe 2+.

rich points analysed in chlorite (XFe= 5 ) were measured in sample SM30 which contains more than 20% of modal haematite. (3) The compositions of the chlorites which sealed the porosity of the phyllic veins in association with magnetite +__phengite correspond to Fe-rich brunsvigite; their mean XF~ values range from 0.54 to 0.68 (Table 8). Chemically they differ from the chlorites associated with the epidotes by higher Mn, higher A1 and lower silicon contents. They also differ by a wide range of intrasample compositional variation (XFe standard deviation ranging from

0.05 to 0.09) which is due to progressive chemical change with respect to the distance from the median part of the veins. This chemical zoning is characterized by a decrease in iron and manganese content away from the vein (Fig. 6 ). Magnetite is restricted to patches of the Fe-Mn-richest chlorite in the median part of veins. Points analysed in chlorite may attain Xve values up to 0.80 in the median part of veins (sample SM45). The composition of phengites coexisting with iron rich chlorites are summarized in Table 9. They are Fe-rich phengites (3 to 4.5 wt.% FeO )

CHEMICALVARIATIONSIN ASSEMBLAGESINCLUDING EPIDOTE AND/OR CHLORITE

105

TABLE 6 Mean microprobe analyses and structural formulas o f chlorites coexisting with epidote, in absence of iron oxide mineral, in volcanic and sedimentary host rocks Samples: An. in aver.:

SM56 27

SM04 17

SM09 19

SM 10 18

SM 14 l0

SM 17 12

SM21 14

SM33 15

SiO2 FeO MgO TiO2 MnO CaO Na20 K20 Total

28.15 17.29 23.02 15.49 0.01 0.43 0.20 0.06 0.07 85.55

27.57 20.03 22.17 15.81 0.13 0.79 0.18 0.02 0.05 86.75

26.69 20.47 24.77 14.47 0.06 0.73 0.09 0.04 0.04 87.35

28.76 17.56 19.58 19.30 0.06 0.40 0.09 85.75

29.00 19.26 22.66 16.66 0.07 0.57 0.09 0.03 0.04 88.38

28.67 19.92 20.31 17.50 0.91 0.13 0.03 86.47

29.83 19.64 15.56 20.28 0.70 0.07 0.05 86.13

28.58 18.07 23.45 17.45 0.04 0.7 l 0.08 0.04 88.42

Si AI Iv AIvx Fe 2÷ Mg Ti Mn 2+ Ca Na K Oct )tEe

6.06+0.14 1.94 2.46+0.21 4.15+0.09 4.99+0.20 0.08 0.05 0.02 0.02 11.77 0.45--+0.02

5.77+0.13 2.23 2.71 +0.15 3.89+0.12 4.93+0.16 0.02 0.14 0.04 0.01 0.01 11.75 0.45+--0.02

5.63+0.10 2.37 2.72+0.15 4.37+0.08 4.55+0.08 0.01 0.13 0.02 0.02 0.01 11.83 0.50+0.01

5.99+0.08 2.01 2.30+0.14 3.41+0.23 5.99+0.23 0.01 0.07 0.02 11.80 0.37+-0.03

5.94+0.06 2.06 2.60+0.11 3.88+0.22 5.09+0.16 0.01 0.10 0.02 0.01 0.01 11.72 0.44-+0.03

5.75+0.28 2.25 2.63+0.10 3.53+0.27 5.42+0.06 0.16 0.03 0.01 11.78 0.41 -+0.02

6.03+0.09 1.97 2.71 +0.09 2.63+0.24 6.11+0.21 0.12 0.02 0.01 11.61 0.31 +-0.03

5.91+0.04 2.09 2.30+0.09 4.06+0.15 5.38+0.08 0.01 0.12 0.02 0.01 11.89 0.44+-0.02

AI203

Structural formulas have been calculated on the basis of 28 oxygens; total iron has been considered as Fe 2+; XFe= 100 (Fe 2+ ) / (Fe 2+ q- Mg 2+ + Mn 2+ ); +- = standard deviation; Oct = octahedral occupancy.

7.0

/ Vein

o

...~

6.0

+ ¢q

•. ' ~°'"% 5.0

including a low but significant content of Mn (0.05 to 0.26 wt.%). The compositional data ofphengites from sample SM45 show low Si and K content and an excess of octahedral occupancy (4.16 atoms); this is probably due to mixing of a few chlorite flakes with the phengitic matrix.

.k;".Farthest wall-rock 0

i

!

i

u

.2

.4

.6

.8

i

u

1.0

1.2

Mn2+ atoms Fig. 6. Atomic iron versus atomic manganese for chlorite associated with phengite + magnetite in a phyllic veinlet e n v i r o n m e n t (sample SM45 ). The solid line (with an arrow) indicates the compositional change o f chlorites with increasing distance f r o m the m e d i a n part o f the vein. The maximal distance o f chlorite occurrence is about 3 c m from the vein.

Miissbauer spectra of chlorites Figure 7 illustrates the M/Sssbauer spectra and Table 10 gives the computed parameters obtained for four chlorite concentrates. All the chlorite spectra show two F e 2+ doublets and one or two Fe 3+ doublets with respective quadrupole splitting, isomer shift and width at half height which agree with the literature data

106

D. BEAUFORTET AL.

TABLE 7 Mean microprobe analyses and structural formulas of chlorites coexisting with epidote + haematite in volcanic and sedimentary host rocks Samples: An. in aver.:

SM23 13

SM26 13

SM27 9

SM29 10

SM30 11

SM19 7

SiO2 A1203 FeO MgO TiO2 MnO CaO Na20 K20 Total

30.38 17.80 13.22 25.15 0.02 0.60 0.13 0.03 0.04 87.37

29.46 19.72 18.04 20.09 0.05 0.69 0.17 0.03 88.25

30.60 20.00 13.54 23.13 0.42 0.05 0.01 0.03 87.78

29.58 19.24 14.77 23.77 0.59 0.10 88.05

32.47 19.68 4.01 31.19 0.05 0.63 0.09 0.04 88.16

31.94 18.94 5.80 28.91 1.76 0.05 87.40

Si AI TM AL vl Fe 2+ Mg Ti Mn 2÷ Ca Na K Oct X~e

6.00 + 0.12 2.00 2.15+0.14 2.20 +_0.16 7.43+0.12 0.10 0.03 0.01 0.01 11.93 0.24 _+0.02

5.90_+ 0.13 2.10 2.56+0.08 3.03 + 0.09 6.00+0.11 0.01 0.12 0.04 . 0.01 11.77 0.34 _+0.01

6.00 + 0.04 2.00 2.62+0.08 2.22 + 0.09 6.76+0.09 0.01 0.07 0.01 .

. 0.01 11.70 0.25 _+0.01

5.85 + 0.12 2.15 2.33+0.13 2.45 + 0.18 7.01 _+0.23 0.10 0.02 . 11.91 0.27 _+0.02

6.05 + 0.05 1.95 2.37+0.05 0.62 + 0.03 8.66_+0.09 0.11 0.02

6.09 _+0.07 1.91 2.34+0.12 0.93 + 0.13 8.21 _+0.20 0.01 0.29 0.01

0.01 11.79 0.08 _+0.01

11.79 0.13 _+0.02

Structural formulas have been calculated on the basis of 28 oxygens; total iron has been considered as Fe2+; XFe= (Fe 2+ ) / (Fe 2÷ + M g 2+ + M n 2+ ); _+ = s t a n d a r d deviation; Oct = octahedral occupancy.

( G o o d m a n and Bain, 1979; Ballet et al., 1985; Borggaard et al., 1982; among others). A weak Fe 3+ doublet with characteristic Mrssbauer parameters of epidote (Dollase, 1973) indicates the presence of minor epidote in the chlorite concentrate of sample SM04. These data show that, in these chlorites, ferrous iron represents 68 to 83% of the total Fe (and ferric iron varies between 16 and 32% of the total Fe ). The higher Fe 3+ content (32%) has been found in chlorites which coexist with haematite and Fe 3+-rich epidotes (SM26). The lower Fe 3+ content (16%) has been found in chlorite which coexist with magnetite and phengites (SM 104 ). The intermediate Fe 3+ contents (30 and 25%) have been found in chlorites associated with epidotes without iron oxide (SM04 and SM56).

Interpretation When one examines the lateral distribution of secondary minerals within the roots of the fossil geothermal system of Saint Martin, epidote a n d / o r chlorite may be continuously observed as far as 3 km from the quartz-diorite pluton. The chemical compositions measured at individual points in epidote and chlorite indicate wide and concomitant variations. They vary with respect to the Ps + Pm ratio from 15 to 36 in the epidotes and with respect to XFe from 0.80 to 0.05 in the chlorites. These chemical variations span the compositional range of epidotes and chlorites observed in most geothermal systems and propylitized area (Bird et al., 1984; Shikazono, 1984; Liou et al., 1985; Shikazono and Kawahata, 1987 ). Considering

CHEMICAL VARIATIONS IN ASSEMBLAGESINCLUDING EPIDOTE AND/OR CHLORITE

S M 26

TABLE 8

T

Mean microprobe analyses and structural formulas of chlorites coexisting with phengite + magnetite in phyllic veins Samples: SM25

"1

z1

o

I--

SM04

!

107

i

An. in aver.;

13

SM43 15

SM40 10

SM45 48

SiO2 A1203 FeO MgO TiO2 MnO CaO N a 20 K20

26.15 20.67 28.74 11.06 0.06 1.77 0.09 0.02 88.56

25.59 19.18 27.03 13.23 0.06 0.92 0.08 0.01 0.03 86.13

26.09 20.56 25.37 12.49 0.06 1.59 0.09 0.05 86.30

25.70 20.00 31.88 8.96 0.06 2.28 0.09 0.01 0.05 89.03

5.62+0.17 2.38 2.84+0.17 4.57+0.31 4.01+0.37 0.01 0.29 0.02 0.01 11.75 0.55+0.05

5.58+0.17 2.42 2.70+0.26 5.79+0.56 2.90+0.82 0.01 0.42 0.02 0.01 11.85 0.68+0.09

Total Si Allv A1vl Fe 2÷ Mg Ti Mn 2÷ Ca Na K

SM56

Oct XF~

5.57+0.25 2.43 2.76+0.24 5.12+0.45 3.51+0.42 0.01 0.32 0.02 0.01 11.75 0.61+ 0.06

5.58+0.14 2.42 2.51 +0.30 4.93+0.26 4.30+0.35 0.01 0.17 0.02 0.01 11.95 0.54+0.05

Structural formulas have been calculated on the basis of 28 oxygens; total iron has been considered as Fe2+; Xv, = (Fe 2+ + Mn 2+ ) / (Fe 2+ + Mg 2+ + Mn 2+ ); + = standard deviation; Oct = octahedral occupancy.

SM45

,~,

~'

~

temperature and oxygen fugacity on the chemical variations of epidote and chlorite.

Phase relations VELOCITY

mm/s

Fig. 7. M6ssbauer spectra o f the studied chlorite. Solid

c u r v e s = d o u b l e t s attributed to Fe 2+ and Fe 3+ in the octahedral or brucitic sites o f chlorites. Dashed curve in sample SM04 corresponds to Fe 3+ in M3 site o f minor epidote.

the existing data on geothermometry (Beaufort et al., 1990) and the information which can be deduced from the local presence of iron oxides (haematite or magnetite) and from Mtissbauer spectra of chlorites, the fossil geothermal system of Saint Martin constitutes a propitious case to evaluate the influence of

Figure 8 gives the diagnostic mineral assemblages in the extensive zone of Saint Martin hydrothermal alteration (illustrated in an A C - F diagram) as a function of the fossil thermal profile. Even if coexisting epidote and chlorite are widely extended in this zone, it can be seen that this association is not always verified: - - Epidote coexists with actinolite in weakly fractured rocks at temperatures higher than 300°C near the periphery of the intrusive body. m Chlorite associated with phengite and magnetite (not shown in Fig. 8) appears without

108

D. BEAUFORT ET AL.

TABLE 9 Mean microprobe analyses and structural formulas of phengites associated with chlorite_+ magnetite in phyllic veins Samples: An. in aver.:

SM25 15

SM40 10

SM45 5

SiO2 A1203 FeO MgO TiO2 MnO CaO Na20 K20 Total

48.84 31.31 3.55 1.26 0.26 0.15 0.02 0.10 10.69 96.18

49.47 30.44 4.47 1.40 0.05 0.07 0.08 10.40 96.37

47.98 33.17 3.28 1.01 0.16 0.20 0.01 0.20 10.31 96.32

Si AITM AIvl Fe 2+ Mg Ti Mn 2÷ Ca Na K Oct XF~

6.52 _+0.05 1.48 3.45 + 0.07 0.31 +0.04 0.25 _+0.04 0.03 0.02 0.03 1.82 + 0.05 4.06 55

6.62 _+0.08 1.38 3.43 ___0.14 0.35_+0.05 0.28 _+0.08 0.01 0.01 0.02 1.78 _+0.08 4.07 56

6.36 _+0.05 1.64 3.54 + 0.07 0.36_+0.05 0.20 _+0.04 0.02 0.02 0.05 1.74 _+0.04 4.16 64

Structural analyses have been calculated on the basis of 22 oxygens; total iron has been considered as Fe 2÷; XF~= (Fe 2+ + M n 2+ ) / ( F e 2+ + M g 2+ + M n 2+ ); + =standard deviation; Oct = octahedral occupancy. TABLE 10 Computer Mrssbauer parameters of purified chlorites of Saint Martin Sample

SM26 SM04 SM56 SM45

Fe 2+

Fe 3+

A

6

F

%

A

6

F

%

2.64 2.33 2.63 2.23 2.62 2.30 2.64 2.30

1.14 1.12 1.14 1.14 1.14 1.11 1.14 1.17

0.27 0.33 0.28 0.29 0.29 0.35 0.31 0.40

50 18 59 11 64 11 57 27

0.70 1.25 0.84

0.37 0.40 0.37

0.49 0.60 0.57

18 14 30

0.69 1.39 0.80

0.39 0.36 0.37

0.44 0.40 0.51

18 7 16

All values in m m s - J. I s o m e r shifts relative to iron metal.

epidote in veins with strongly phengitized wall rocks. The transition chlorite-actinolite occurred near 300 °C (SM33) and the transition mixed-

layer chlorite/smectite-chlorite near 200220°C. Epidote and chlorite coexist between 200 and 300 ° C. All these thermal estimations are consistent with temperatures reported for similar mineral transitions in most other geothermal environments (Ellis, 1979; Bird et al., 1984; Kristmannsdottir, 1979; Franzson et al., 1986). At temperatures lower than 200°C, calc-silicate minerals did not appear but calcite crystallized with chlorite/smectite and illite/smectite mixed-layer minerals were abundant. The formation temperature of chlorites in strongly phengitized veins is difficult to estimate because no fluid inclusions may be clearly related to the chlorite crystallization stage. They are inferred to have crystallized below 250°C because they sealed the residual porosity between early quartz crystals for which study of fluid inclusions indicated temperatures varying between 330 and 250 ° C.

Influence of temperature on the chemical variations of epidote and chlorite Even if we have seen in Figure 8 that temperature exerted a major influence on the transition of silicate phases and consequently controlled the phases equilibrium, its influence on the compositional variations observed in epidotes and chlorites of the extensive hydrothermal area is not evident. The variation of Ps content of epidote and the variation of XFe in chlorite cannot be satisfactorily correlated with the temperatures. In this respect, the use of the various geothermometers based on the compositional variations of chlorites is not valid in this environment. The most interesting chemical variations which can be correlated with the paleotemperature profile is the standard deviation of the Ps + Pm content of epidote and to a lesser degree the standard deviation of the aluminium content of chlorite (Fig. 9 ). It can be seen that the scatter of composition ofepidote and chlorite substantially increases at low temperatures

109

CHEMICAL VARIATIONS IN ASSEMBLAGES INCLUDING EPIDOTE AND/OR CHLORITE A

/ / A

c

C /

\

\ \

~, ~b . act

~'F'

(°c) ~l;n . ~u

"/>", j

20o

~

;

o

DISTANCE FROM THE INTRUSION (kin) Fig. 8. Phase relationships of the mineral assemblages in the extensive zone of fossil hydrothermal alteration at Saint Martin, illustrated in terms of A - C - F diagrams, as functions of paleotemperature and distance from the thermal source (quartz-diorite pluton). The chlorite + phengite + magnetite assemblage observed in veins is not reported in this figure. ep = epidote; act = actinolite; chl= chlorite; mag= magnetite; hm = haematite; cal= calcite; c/s = chlorite/smectite mixedlayer mineral; i/s= illite/smectite mixed-layer mineral.

(220-240 oC ). Such a trend reveals an increasing deviation of the chemical equilibrium in both epidotes and chlorites with the decrease of their formation temperature away from the thermal source. A similar trend has been already reported for chlorite in the Salton Sea geothermal field by Cho et al. ( 1988 ). Velde et al. ( 1991 ) described a similar compositional evolution for low-temperature chlorites in diagenetic environments and suggested that these minerals crystallized with varying compositions and then come into chemical equilibrium by the combined influence of temperature and post-crystallization time on solid state diffusion processes. Patrier et al. ( 1991 ), in a study of Fe3+/A1 substitutional order/disorder in epidotes from the geothermal system of Saint Martin, demonstrated that the crystals

which formed below 300°C (in association with chlorite) exhibit a metastable ordering state which clearly increases toward the outer zone of the system in which temperatures were low. These authors interpreted this variation of ordering state in epidotes as the result of the "annealing" of disordered original crystals, during the post-crystallization geothermal history, by solid state diffusion processes. It is particularly interesting to note that this variation coincides with the variation of chemical scatter presented in this work. All these considerations argue that the epidotes and chlorites coprecipitated at Saint Martin must be considered as phases of which the present chemical metastability increases with distance from the intrusion. This appears in response to the decrease of temperature and

1 10

D. BEAUFORT ET AL.

Ps+Pm in epidotes 5 _ 4

3

2

Z 0

•

OU

• O 0

1

-

-

-

- 0 -

_

•

P 0

!

~,4o

200 AI atoms in chlorites

~e

.<

Z

.20

'

:~eo

'

~20

'

TEMPERATURE

o~ o~

.10

~e.

•

•

b- o-. . . . . . O0

T

200

~+o

'

• "-'O

.......

~8o

-O

'

~2o

'

TEMPERATURE

Fig. 9. Plot of standard deviations of mean Ps + Pm content of epidotes and standard deviation of mean atomic AI content of chlorites versus temperature, in the weakly fractured rocks from the fossil system of Saint Martin.

AI

of the time during which these temperatures affected the rocks.

Influence of oxygen fugacity on the chemical variations of epidote and chlorite

+

Mg

I °p

Fe*+ M n

Fig. 10. Plot of assemblages bearing epidote and/or chlorite in A I - M g - ( F e * + M n ) coordinates. Fe*=total Fe; ep=epidote; act=actinolite; chl=chlorite; mag= magnetite; hm = haematite; ph = phengite.

Variations of oxygen fugacity have been indicated by the local presence of different iron oxides and by changes in their nature (haematite or magnetite). The local increasing of oxygen fugacity in the host-rock formations controlled haematite crystallization and several concomitant chemical trends in both epidote and chlorite: - - increasing of Ps content of epidote up to 3032; - - decreasing of XFe content in chlorite down to 0.08; - - increasing of Fe 3+/Fe 2÷ ratio in chlorite up to 32%.

CHEMICALVARIATIONSIN ASSEMBLAGESINCLUDINGEPIDOTEAND/OR CHLORITE

111

Fe3 + + Mn3 + in epidotes

2.00

SM26 ~ / /

1.80

/ /

.O

/

J

/ /

1.60

/"

SM04

/ I / / / /

t

/

1.40

/

SM56

/

1.20

1.00 20

Fe3+/Fe 2+ + Fe3+ in chlorites Fig. 11. Plot of mean Fe 3+ + M n 3+ determined by microprobe in epidotes versus F e 3 + / ( F e 2+ + F e 3+ ) determined in coprecipitated chlorites by M6ssbauer spectrometry in samples SM26, SM56 and SM04.

In Figure 10, we have plotted in an AI-MgFe*+ Mn diagram the mean chemical composition of each silicate which constitutes the different assemblages including epidote a n d / o r chlorite in the hydrothermal system of Saint Martin. In this diagram, the compositional fields of epidotes and chlorites are clearly individualized as a function of their coexisting iron-bearing phase: epidote Ps = 21-22 coexists with actinolite and magnetite near the intrusion; - - all the chlorites which coexists with epidotes are XF~< 0.50; - - e p i d o t e P s = 2 1 - 3 0 coexists with chlorite XF~= 31--50 in absence of iron oxide; epidote P s - 3 0 - 3 2 coexists with chlorite XF~= 8--34 in presence of haematite; chlorite XF~= 54--68 coexist with magnetite and phenglte in veins. It is well known from experimental studies that fo~ strongly controls the Ps content of epidotes; Holdaway (1972) and Liou (1973) demonstrated that epidote is most iron-rich

(Ps= 33 _+2) in the presence of a haematite oxygen buffer and becomes more aluminous with decreasing fo~. Whereas the XFe content of chlorites is highly related to the XFe content of bulk rock in most of the studied hydrotherreal area, it appears that this assumption is not valid in haematite-bearing rocks for which XFe contents in bulk rock and in chlorite are negatively correlated (Fig. 5b). This fact is a consequence of an expected structural limitation in the assignment of Fe 3+ in trioctahedral chlorite which favours the incorporation of magnesium (rather than ferric iron) in chlorite under oxidizing conditions as demonstrated by Rumble (1976) and Bryndzia and Scott (1987). A compilation of Mrssbauer spectra of chlorites published in the literature (see references cited above in "results of MOssbauer spectra of chlorites") reveals that the Fe 3+/Fe 2+ ratio in chlorites rarely exceed 30% in natural systems. Despite a small number of analyses, the Fe 3+/ Fe 2+ ratio measured in chlorites by M~issbauer

112

spectrometry in this work corroborates this fact and points to a good relationship between measured Fe3+/Fe 2÷ ratio in chlorites and conditions of oxygen fugacity expected from coprecipitated iron oxides. Chlorites with a Fe 3+/Fe 2÷ ratio of 32O/ocrystallized with haematite at nearly the same temperature as chlorites with a Fe3+/Fe 2÷ ratio of 16% crystallized with magnetite. Figure 1 1 shows a distribution of Fe 3+ between chlorites and epidotes and reveals that the strong compositional changes observed between epidotes of samples SM04 and SM56 in the outer part of the epidote-chlorite zone were controlled by different conditions of oxygen fugacity. In summary, the significant compositional variations of epidotes (Ps content) and of chlorites (XFe) are the result offo2 variations in the extensive hydrothermal system of Saint Martin. With increasingfo2 in solutions (up to the haematite buffer), the Fe 3÷ enrichment of epidote is concomitant with Mg enrichment of coexisting chlorite. The causes of fo~ were probably multiple: local influence of bulk rock chemistry controlling haematite crystallization; - - composition of infiltrated fluids in veins environment; decreasing of oxygen fugacity by thermal metamorphism near the intrusion. Conclusions

The compositional variations of epidote a n d / o r chlorite of four distinct assemblages in the extensive hydrothermal zone of the geothermal system of Saint Martin are strongly influenced by the combined effects of temperature and fo:- At Saint Martin, the effect offo: is particularly perceptible in the control of epidote Ps content, of chlorite XFe ratio and of Fe 3+ distribution between coexisting epidotes and chlorites. Even if it is well demonstrated by changes in phase relations, the effect of temperature is not so clear at the scale of compositional ranges because epidote and chlorite

D. BEAUFORT ET AL.

have not reached the chemical equilibrium. Despite the fact that it may be partially canceled out by the effect offo2, the effect of temperature can be found in the variation of compositional ranges of both phases which increase toward the outer part of the geothermal system in response to the combination of decreasing temperatures and decreasing time of heating of the rocks. However, in deep geothermal systems such as that of Saint Martin, temperature and fo2 were not always independent controlling factors. This was probably true near the periphery of the intrusion: with increasing metamorphic grade, the decrease of oxidation state of rocks was concomitant with increasing temperature. This study rises the problem of the validity of purely thermodynamical approaches frequently used for the determination of mineral geothermometers in hydrothermal environments. The use of these models based on ideal chemical equilibrium between phases in systems which have attained the phase equilibrium but not the chemical equilibrium may lead to inaccurate or simply incorrect results. Acknowledgments

Financial support of this study was provided by the "Programme interdisciplinaire de recherche sur les sciences pour l'6nergie et les mati6res premi6res". References

Ballet, O., Coey, J.M.D. and Burke, K.J., 1985. Magnetic propertiesof sheetsilicates;2:1 : 1 layerminerals.Phys. Chem. Miner., 12: 370-378. Beane, R.E., 1982. Hydrothermal alteration in silicate rocks, southwestern North America. In: S.R. Titley (Editor), Advancesin Geologyof the Porphyry Copper Deposits,SouthwesternNorth America.Univ. Ariz. Press, AZ, Tucson, Chapter 6, pp. 117-137. Beaufort, D. and Meunier, A., 1983.A petrographicstudy of phyllic alteration superimposedon potassic alteration: The Sibert porphyry deposit (Rhone, France). Econ. Geol., 78: 1514-1527. Beaufort, D., Westercamp, D., Legendre, O. and Meu-

CHEMICALVARIATIONSIN ASSEMBLAGESINCLUDINGEPIDOTEAND/ORCHLORITE nier, A., 1990. The fossil hydrothermal system of Saint Martin: ( 1 ) Geology and lateral distribution of alterations. J. Volcanol. Geotherm. Res., 40:219-243. Bettison, L.A. and Shiffman, P., 1988. Compositional and structural variations of phyllosilicates from the Point Sal ophiolite, California. Am. Mineral., 73: 62-76. Bird, D.K. and Helgeson, H.C., 1980. Chemical interaction of chemical solutions with epidote-feldspar mineral assemblages in geological systems. 1. Thermodynamic analysis of phase relations in the system CaOFeO-Fe203-A1203-SiO2-H20-CO2. Am. J. Sci., 280: 907-941. Bird, D.K. and Helgeson, H.C., 1981. Chemical interaction of chemical solutions with epidote-feldspar mineral assemblages in geological systems. 2. Equilibrium constraints in metamorphic/geothermal processes. Am. J. Sci., 281: 576-614. Bird, D.K., Shiffman, P., Elders, W.A., Williams and Mac Dowell, S.D., 1984. Calc-silicate mineralization of active geothermal systems. Econ. Geol., 79:671-695. Bird, D.K., Cho, M., Janik, C.J., Liou, J.G. and Caruso, L.J., 1988. Compositional, order/disorder, and stable isotope characteristics of AI-Fe epidote, State 2-14 drill hole, Salton Sea geothermal system. J. Geophys. Res., 93: 13,135-13,144. Borggaard, O.K., Lindgreen, H.B. and Morup, S., 1982. Oxidation and reduction of structural iron in chlorite at 480°C. Clays Clay Miner., 5: 353-364. Bryndzia, L.T. and Scott, S.D., 1987. The composition of chlorite as a function of sulfur and oxygen fugacity: An experimental study. Am. J. Sci., 287: 50-76. Caruso, L.J., Bird, D.K., Cho, M. and Liou, J.G., 1988. Epidote-bearing veins in the State 2-14 drill hole: Implications for hydrothermal fluid composition. J. Geophys. Res., 93: 13,123-13,133. Cathelineau, M. and Nieva, D., 1985. A chlorite solid solution geothermometer. The Los Azufres (Mexico) geothermal system. Contrib. Mineral. Petrol., 91: 235244. Cathelineau, M., Oliver, R., Nieva, D. and Garfias, 1985. Mineralogy and distribution of hydrothermal zones in Los Azufres (Mexico) Geothermal field. Geothermics, 14: 49-57. Cavaretta, G., Fianelli, G. and Puxxedu, M., 1982. Formation of Authigenetic minerals and their use as indicators of the physiochemical parameters of the fluid in the Larderello-Travale geothermal field. Econ. Geol., 77: 1071-1084. Cho, M., Liou, L.G. and Bird, D.K., 1988. Prograde phase relations in the state 2-14 well metasandstones, Salton Sea Geothermal field, California. J. Geophys. Res., 93: 13,081-13,113. Creasey, S.C., 1959. Some phase relations in hydrothermally altered rock of porphyry copper deposits. Econ. Geol., 54:351-373. Deer, W.A., Howie, R.A. and Zussman, M.A., 1962. Rock

113

Forming Minerals. Vol. 1; Ortho and Ring Silicates. Longman, New York, NY, 333 pp. Dollase, W.A., 1973. M6ssbauer spectra and iron distribution in the epidote group minerals. Z. Kristallogr., 138: 41-63. Elders, W.A., Hoagland, J.R., Mac Dowell, S.D. and Corbo, J.M., 1979. Hydrothermal mineral zones in the geothermal reservoir of Cerro Prieto, Mexico. Geothermics, 8: 201-209. Ellis, A.J., 1979. Explored geothermal systems. In: H.L. Barnes (Editor), Geochemistry of Hydrothermal Ore Deposits. Wiley, New York, NY, 2nd ed., pp. 305-360. Exley, R.A., 1982. Electron microprobe studies of Iceland research drilling project high-temperature hydrothermal mineral geochemistry. J. Geophys. Res., 87: 65476557. Fawcett, J.J. and Yoder, H.S., 1966. Phase relationships of chlorites in the system MgO-A1203-SiO2-H20. Am. Mineral., 51: 353-380. Fleming, P.D. and Fawcett, J.J., 1976. Upper stability of chlorite + quartz in the system igO-FeO-Al203-SiO2H20 at 2kbars water pressure. Am. Mineral., 61:11751193. Foster, M.D., 1962. Interpretation of the composition and a classification of the chlorites. U.S. Geol. Surv., Prof. Pap., 414-A: 1-33. Franzson, H., Gudmundsson, A., Fridleifsson, G.O. and Tomasson, J., 1986. Nesjavellir high-t field, SW Iceland-reservoir geology. Proc. Fifth Int. Symp. WaterRock Interaction, pp. 210-213. Goodman, B.A. and Bain, D.C., 1979. M6ssbauer spectra of chlorites and their decomposition products. In: Proc. Int. Clay Conf., Oxford, 1978, M.M. Mortland and V.C. Farmer (Editors), Elsevier, Amsterdam, pp. 65-74. Hietanen, A., 1974. Amphibole pairs, epidote minerals, chlorite and plagioclase in metamorphic rocks, Northern Sierra Nevada, California. Am. Mineral., 59: 2240. Holdaway, M.J., 1972. Thermal stability ofA1-Fe epidote as a function offo: and Fe content. Contrib. Mineral. Petrol., 37: 307-340. Huelen, J.B. and Nielson, D.L., 1986. Hydrothermal alteration in the Baca geothermal system, Redondo dome, Valles caldera, New Mexico. J. Geophys. Res., 91: 1867-1886. Joesten, R. and Fisher, G., 1988. Kinetics of diffusion controlled mineral growth in the Christmas Montains (Texas) contact aureole. Geol. Soc. Am. Bull., 100: 714-732. Kristmannsdottir, H., 1979. Alteration of basaltic rocks by hydrothermal activity at 100-300°C. In: M.M. Mortland and V.C. Farmer (Editors), Int. Clay Conf. 1978. Elsevier, Amsterdam, pp. 359-367. Laird, J., 1988. Chlorites: Metamorphic petrology. In: S.W. Bailey (Editor), Hydrous PhyUosilicates (exclusive of micas). Reviews in mineralogy, 19, P.H. Ribbe series, Mineralogical Society of America, pp. 405-447.

114 Liou, J.G., Hyung, Shik KJm and Maruyama, S., 1983. Prehnite-epidote equilibria and their petrologic application. J. Petrol.,2 4: 321-342. Liou, J.G., Seki, Y., Guillemette, R.O. and Seki, H., 1985. Compositions and parageneses of secondary minerals in the Okinabe geothermal system, Japan. Chem. Geol., 49: 1-20. Liou, J.L., 1973. Synthesis and stability relations of epidote, Ca2AlzFeSi30~2(OH). J. Petrol., 14:381-413. Lowell, J.D. and Guilbert, J.M., 1970. Lateral and vertical alteration and mineralization zoning in porphyry ore deposits. Econ. Geol., 65: 373-408. Mac Dowell, S.D. and Elders, W.A., 1980. Authigenetic layer silicate minerals in Borehole Elmore 1, Salton Sea geothermal field, California, U.S.A. Contrib. Mineral. Petrol., 74:293-3 I0. Mac Onie, A.W., Fawcett, J.J. and James, R.S., 1975. The Stability of intermediate chlorite of the clinochloredaphnite series at 2Kb PH20. Am. Mineral., 60: 10471063. Myashiro, A. and Seki, Y., 1958. Enlargement of the compositional field to epidote and piemontite with rising temperature. Am. J. Sci., 256: 423-430. Nelson, B.W. and Roy, R., 1958. Synthesis of chlorites and their structural and chemical constitution. Am. Mineral., 43: 707-725. Patrier, P., Beaufort, D., Meunier, A., Eymery, J.P. and Petit, S., 1991. Determination of the nonequilibrium ordering state in epidote from the ancient geothermal field of Saint Martin: application of Mt~ssbauer spectroscopy. Am. Mineral., 76: 602-610.

D. BEAUFORTETAL. Rumble, D., 1976. Oxide minerals in metamorphic rocks. In: D. Rumble III. (Editor), Oxide Minerals. Mineral. Soc. Am., Short Course Notes, 3: R1-R24. Shikazono, N., 1984. Compositional variations in epidotes from geothermal areas. Geochem. J., 18: 181187. Shikazono, N. and Kawahata, H., 1987. Compositional differences in chlorite from hydrothermally altered rocks and ore deposits. Strens, R.G.J., 1965. Stability relations of the AI-Fe epidotes. Mineral. Mag., 35: 464-475. Stoessel, R.K., 1984. Regular solution site-mixing model for chlorites. Clays Clay Miner., 32: 205-212. Thomasson, J. and Kristmannsdottir, H., 1972. High temperature alteration minerals and thermal brines, Reykjanis, Iceland. Contrib. Mineral. Petrol., 36: 123134, Velde, B., 1973. Phase equilibria study in the system MgOAI2Oa-SiO2-H20: Chlorites and associated minerals. Mineral. Mag., 39:297-312. Velde, B. and Rumble, D., III., 1977. Alumina content of chlorite in muscovite bearing assemblages. Geophys. Lab. Annu. Rep., 1976-1977: 621-623. Velse, B., El Moutaouakkil, N. and Iijima, A., 1991. Compositional homogeneity in low-temperature chlorites. Contrib. Mineral. Petrol., 107: 21-26. Walshe, J.L., 1986. A six component Chlorite solid solution model and the conditions of chlorite formation in hydrothermal and geothermal systems. Econ. Geol., 81: 681-703.